Diversity of volcanism and evidence of discrete eruption centres in the Miocene Andes of Central Chile

doi:10.21203/rs.3.rs-3016212/v1

Download PDF

Research Article

Diversity of volcanism and evidence of discrete eruption centres in the Miocene Andes of Central Chile

https://doi.org/10.21203/rs.3.rs-3016212/v1

This work is licensed under a CC BY 4.0 License

Version 1

posted

You are reading this latest preprint version

The geology of the Farellones region in the Andean Cordillera of central Chile, comprise a thick sequence of volcanic and volcaniclastic Miocene rocks. The occurrence of discrete eruption centres within this sequence, whilst suggested, has been difficult to stablish, as in the modern volcanic systems of the SVZ to the southeast of the study area. Here we report, for the first time a previously uncharacterized composite (Cerro Colorado Volcano) associated to a series of intercalated lava flows and pyroclastic deposits of basaltic to rhyolitic composition of Miocene age. Eruptive sequences reach 1.7 km in vertical section, yet the lateral continuity of units is interrupted by a major faults, and gravity collapse deposits. We document a series of large pyroclastic block and ash flows, which are overlain by rhyolitic pyroclastic density currents, ash fall and crystal, lithic and vitric tuffs, all making up more than 300 m of the sequence. A further 300 m thick unit of spherulite-bearing rhyolitic lavas represent the upper most section of the Cerro Colorado volcano. During the early stages of its construction magmas and hydrothermal fluids interacted to form a maar-diatreme system (Quebrada Lunes maar-diatreme). Much of the volcanic units are intruded by E-W and NW-SE striking andesitic and rhyolitic dikes One of the dikes exhibits internal pyroclastic textures evidence of shallow conduit fragmentation. This pyroclastic dike is further evidence of magma-fluid interaction and related explosivity. Similar pyroclastic breccias described elsewhere related to porphyry copper mineralization. Our results indicate that Miocene volcanism in the Andes of Central Chile is linked to discrete eruptive centres which can be identified and characterized through careful field mapping.

Miocene volcanism

maar-diatreme

rhyolitic domes

pyroclastic dike

lava flow

explosive volcanic deposits

Andes of Central Chile

Miocene volcanic and volcaniclastic rocks, in the western Andean Cordillera to the east of Santiago in Central Chile (Fig. 1), have been generally included in a single volcanic unit (Farellones Formation Beccar 1983; Vergara et al. 1988; Rivano et al. 1990; Thiele et al. 1991; Nyström et al. 2003; Muñoz et al. 2006)), with unconformably overlies the volcanic Eocene-Oligocene Abanico Formation. Even if several studies describe the petrology, geochemistry and geochronology of the two no studies has been carried out in order to locate and characterize individual volcanic centres within the volcanic sequences. Delimiting ancient Miocene volcanoes in the Andes of central Chile is important since identification of their eruptive products and relative timings may be useful for constraining hazards at active volcanoes nearby and since their roots represent sites of important ore deposits (Serrano et al. 1998).

[FIG. 1]

Mentions to Miocene eruptive centres can be found in (1) Thomas (1953), who stated that the lavas and tuffs of Farellones came from discrete volcanic centres and recognized a fissure vent.; (2) Beccar (1983) indicated that the La Parva, Colorado, and El Pintor hills are sections of an eroded Miocene volcanic centre; (3) Rivano (1990) also suggested that the peaks of “La Parva” and “El Pintor” were a deeply eroded stratovolcano; and finally (4) Thiele et al. (1991) described the Farellones Formation as a sequence of lavas and ash deposits with the uppermost member being a series of domes, laccoliths, considered to be remnants of stratovolcanoes. Later works (e.g., Nyström et al. 2003; Muñoz-Gómez et al. 2006; Piquer et al. 2010; Piquer 2015) focused on the geochemistry and structural characterization but an analysis of physical volcanology in the area has been lacking.

In this contribution we provide new data from geological field mapping and analysis on both the location and characteristics of previously unrecognized volcanoes, that gave rise to volcanism during the Miocene in the Farellones area. These data provide an improved understanding of the volcano eruption style with physical volcanology interpretation and of both constructive and destructive volcanic processes. Through the analysis of high-resolution geological mapping, we were able to discover and describe two unrecognised eruptive centres, describe the predominant eruptive style and determine how the volcanoes were built and destroyed over time. These results will help to improve the understanding of the behaviour of the Miocene volcanic arc in the Southern Central Andes.

In the study area, basaltic to andesitic volcanic rocks related to an Eocene-Oligocene extensional intra-arc basin (Fig. 1) have been described (e.g., Charrier et al. 2002; Piquer et al. 2010). Many of these volcanic rocks became intercalated with lacustrine deposits from the basin itself (Nyström et al. 2003; Mosolf 2013). At the latitudes 33°S to 34°S, this syn-extension Eocene-Oligocene basin deposits are referred to as the Abanico Formation (e.g., Muñoz 1953 in Klohn 1960; Aguirre 1960; Charrier et al. 2002; Nyström et al. 2003; Muñoz-Gómez et al. 2006; Godoy, et al. 1999; Jara & Charrier 2014). The geological formation has been characterized as a > 3.3 km thick sequence divided into two principal members, the lower: formed by dacitic to rhyolitic ash flows, interbedded with lacustrine deposits; and the upper: formed predominantly by basaltic to andesitic lavas (Nyström et al. 2003; Muñoz-Gómez et al. 2006).

During the Early Miocene, the intra-arc basin became tectonically inverted, which led to the deformation of the Eocene-Oligocene basin rocks (Pilger 1984; Thiele et al. 1991; Nyström et al 2003, Fock et al. 2006). This deformation generated asymmetric folds with large (~ 10 km) wavelengths, as well as reactivation of normal faults and nucleation of new thrusts (Giambiagi et al. 2003; Fock et al. 2006; Pinto et al. 2010; Jara et al. 2015, 2018). Synchronous to the inversion, new Miocene volcanic activity began to form and produce further volcanic deposits that became interbedded with lacustrine deposits (Nyström et al. 2003). These volcaniclastic rocks were deposited in a basin on top of deformed Eocene-Oligocene rocks, generating both concordant and discordant, fault-related, contacts (Thiele et al. 1991; Fock et al. 2006; Godoy 2012). These syn-inversion Miocene rocks can be traced along a N-S belt from 31°30’S to 34°45’S, and for this reason the unit is combined, and the rocks are assigned to the ‘Farellones Formation’ (Aguirre 1960).

This unit is characterized by 2.5 km thick of lava flows, lava breccias, pyroclastic deposits and domes with basaltic to rhyolitic compositions (Nyström et al. 2003). Rivano et al. (1990) divided the formation into three main members; the lower member, related with predominantly pyroclastic rocks with interbedded layers of lacustrine sediments, the middle member composed predominantly by a sequence of andesitic-basaltic lavas, and the upper member formed by both mafic and felsic lavas and domes. Some thick ash flow deposits from the base, cut locally by N-S faults, were described and interpreted as caldera related ignimbrites (Rivano 1990; Thiele et al. 1991; Elgueta et al. 1999; Nyström et al. 2003). In the field, the previously defined contact between the Abanico and Farellones formations can be discordant, fault-related or concordant (Thiele et al. 1991; Fock et al. 2006). These Miocene rocks, has been described to be little to no deformed by the continuous inversion of the basin, forming kilometre-long wavelength gentle folds with an amplitude of about 500 m to being mostly horizontal in the uppermost sections (Vergara et al. 1988; Nyström et al. 2003). The main areas of focus in this work are now correlated with the rocks of the middle and upper member of the Farellones Formation (Fig. 2).

[FIG.2]

Due to the similarities between the rocks from the named Abanico and Farellones formations in origin, nature and localization, the differences that set apart these two formations has been discussed in terms of geochronology, geochemistry signature and deformation. For example, between the Abanico and Farellones formations rocks there is a hiatus of about 1 Ma (Piquer 2015). In terms of geochemistry, for instance, Abanico rocks have an SiO₂ average of 1.2 wt% more than the Farellones rocks (Piquer 2015). Furthermore, and perhaps the most compelling distinction, is that the rocks from the Abanico Formation are highly deformed due to the inversion of the basin after the deposition, whereas the Miocene rocks from the Farellones Formation were deposited synchronically with the inversion and are little to no deformed (Charrier et al. 2002; Nyström et al. 2003; Charrier et al. 2015). Such processes are evidenced by the systematic change in Sr-Nd isotope signatures and REE patterns, suggesting an increase in the magma source depth with time, as reflected, for example, in increasing La/Yb ratios (Nyström et al. 2003)

Miocene volcanic and volcaniclastic rocks in the northern part of the study area are intruded by plutonic rocks. The main intrusive bodies in the area are the Rio Blanco – San Francisco batholith, defined by 14 lithofacies between granodiorites and diorites with U-Pb ages in zircon from ~ 20 Ma to 8.16 Ma (Deckart et al. 2005), and the Yerba Loca stock, which is composed of gabbros to monzodiorites, with U-Pb ages in zircon of 14.9 Ma (Deckart et al. 2010). Additionally, Piquer et al. (2015), describes dacitic to rhyolitic porphyry bodies of late Miocene to early Pliocene U-Pb ages in zircon (7.12–4.69 Ma), intruding previous rocks in the zone and carrying Cu-Mo mineralization.

Later, Pliocene-Quaternary volcanism formed the active volcanic arc expressed by the Tupungato (Diaz 1961), Tupungatito (Bertín & Silva 2015), San José (González-Ferrán 1994), Marmolejo and Maipo volcanoes in the Andes of central Chile (Fig. 1). These volcanoes form the Northern section of the Southern Volcanic Zone (NSVZ; Stern & Skewes 1995) and feature a diversity of volcanic activity. The Maipo Volcanic Complex (Sruoga et al. 2005), resides in a collapse caldera with a large rhyolitic pyroclastic density current forming the Pudahuel ignimbrite emission event at ~ 0.45 Ma (Pineda et al. 2021).

With the objective of understanding and characterizing the physical volcanology of the Farellones region, we incorporated new high-resolution (1:10.000 scale) field geological mapping in the Santiago Cordillera, between 376.000 to 383.000 m E and 6.306.500 m to 6.311.500 m N, in an approximately 42 km² area, which includes part of the localities of Farellones, La Parva, Cerro Colorado, Valle Nevado and the Yerba Loca National Park (Fig. 1c). An even higher-resolution (1:5.000 scale) mapping of volcanic deposits was performed focused on four specific areas, determined by the 1:10.000 geologic mapping, which were variably representative of parts of eroded volcanic centres. The mapping was complemented by drone (DJI Air2) pictures and the microscopic description of 18 petrographic samples (A.1) in order to find representative evidence of interpreted volcanic processes.

In the outcrops where volcanic lithofacies present clast supported textures, we performed a clast analysis which allowed comparison of the distribution of clast size, clast roundness and sphericity, as well as thickness of the deposits and composition when possible. Then, we associated these characteristics with a type of volcanic deposit and an interpretation of the origin and associated eruptive style, to link the lithofacies with the presence (or not) of extinct eruptive centres. The clast size analysis was done through a schematization of high-resolution images in the software Illustrator and the measurement of the long (a) and short (b) 2D axis in the software ImageJ to determine the maximum, minimum and average size of each axis. Later these sizes and the b/a ratios were analysed to study if the clasts tended to be more spheric or more elongated. The main limitation of the method is the two-dimensional nature of the analysis, since the clasts could not be measured along their third dimensions, into the outcrops. However, there is valuable insight to be gained from the 2D analysis (Michol et al. 2008), liable to measurement errors.

Regarding the volcanic products, we made an estimation of their volumes, considering the area mapped and calculated with the software ArcGIS, multiplied by the thickness estimated in the field or from drone pictures. This is an approximate estimate for the minimum volume, with simplified geometries (Lefebvre et al. 2013) and thus, susceptible to measurement error.

From the 1:10.000 mapping we identified the general sequence of units, starting at the bottom of the San Francisco gorge, with reworked Oligocene-Eocene volcanoclastic and pyroclastic deposits, which were dipping steeply and weathered. Up in the sequence, from the bottom of the Molina River we encounter a roughly 700 m thick sequence of Miocene basaltic-andesitic lavas intercalated with dacitic to rhyolitic pyroclastic deposits. The upper most mapped units begin at the Farellones village, and continues to the Cerro Colorado hilltop, consisting of chaotically flow-banded rhyolitic lavas, accompanied with sequences of pyroclastic deposits, forming a discrete topographic high. The entire sequence is intruded by a large gabbroic to monzodioritic plutonic body in the NW extreme of the study zone, and there are E-W and NW-SE andesitic and rhyolitic dikes throughout the sequence found around Cerro Colorado. Finally, Quaternary deposits such as moraines and active colluvial and alluvial cones drape about 50% of the study area, however we do not include these deposits directly in our presented maps since our focus is on volcanic related deposits.

From the 1:10.000 general observations we determined four specific areas to perform the 1:5.000 mapping, as these presented unconformities (such as faults, pyroclastic dikes, rocks unseen in the rest of the area, chaotic disposition, and well exposed sections) that we hypothesized may be linked to eruption centres and their volcanic characterization.

4.1. Evidence of maar-diatreme volcanism in Quebrada Lunes Valley

The Quebrada Lunes Valley (Fig. 2) is an almost 1 km wide, in its upper section, and 1.4 km long gorge located at 377.808 m E, 6.310.667 m N (WGS84 UTM 19S). The valley trends westward and exhibits an inverted cone-like morphology characterized by sharp inclined cliffs, active alluvial deposits and multiple streams that join at the lower-most section of the valley. Nine different principal lithofacies outcrop within the valley, as follows (Fig. 3):

PB. Pyroclastic breccia, ~ 500 m thick, no clear dipping, with angular and high sphericity lithic and juvenile clasts that range in size from 1 cm to 35 cm, in a light grey ash rich matrix. The rock is slightly competent, matrix supported, polimict, unsorted and massive; and locally presents hydrothermal alteration.

AdL1. Andesitic to dacitic aphanitic lava, ~ 50 m thick, dips 35°S, with a late-stage alteration that filled some cavities with quartz geodes. This unit is intruded by a sub-horizontal leucocratic sill (LS).

LaT1. Laminated crystal ash tuff, ~ 10 m thick, dips 30°S with inner yellowish grey layers of 10–30 cm thick.

AdL2. Andesitic to dacitic porphyritic plagioclase-bearing lava, ~ 40 m thick, no clear dipping, with grains between 0.5–1 mm in a grey groundmass. The unit is characterized by the presence of a brecciated lava of the same composition along the roof and bottom and is observed to be intruded with a sub-vertical aphanitic dike (AD). This rock has an Fe-oxide patina, and the breccia parts also present a late alteration that filled some cavities with up to 5 cm quartz geodes. This lava is subhorizontal in the northern part of the valley but dipping 25°S in the southern part of the valley.

LaT2. Laminated (10–30 cm) grey ash tuff, ~ 40 m thick, dips 5°S. The outcrop is unweathered with a light grey colour.

PadB. Matrix-supported polimict andesitic pyroclastic breccia, ~ 60 m thick, dips sub-horizontally, with fractured plagioclase microlites (< 0.1 mm) and fractured lithic and juvenile clasts (0.2 mm–5 cm), which are sub-angular and with medium sphericity, contained within an ash rich matrix. The outcrop is altered by Fe-oxides, giving the rocks a dark red colour.

PT. Pyroclastic tuff, ~ 40 m thick, dips around 10°S, with black aphanitic and andesitic porphyritic lithic and juvenile sub-rounded and low sphericity clasts (2 mm–1 cm) and plagioclase phenocrystals (< 1 mm–6 cm), in an ash rich matrix. This rock is oligomictic, matrix supported, unweathered and with a light grey colour.

ApB. Matrix-supported polimict andesitic to dacitic pyroclastic breccia, ~ 30 m thick, no clear dipping, with clasts up to 10 cm of similar aspect ratios to those described in unit PadB. The outcrop is altered by Fe-oxides, giving the rocks a dark red colour. From this layer, a thin section was made of the matrix of this unit, where we found that the matrix is classified as a lithic vitric tuff with juvenile angular, low sphericity fragments (< 1 mm), including volcanic glass fragments, also known as shards (< 3 mm); lithic angular, low sphericity fragments (< 4mm) of andesites and vitric tuffs strongly fractured, and minor plagioclase and quartz fractured crystals (> 2 mm).

LdT. Finely laminated dacitic tuffs, ~ 1 m thick, dips sub-horizontally, with quartz and amphibole microlites, internally laminated about ~ 2 cm. These facies are weathered and presents a light grey colour. The unit marks the upper-most part of the Quebrada Lunes Valley and at this point is overlain by a series of Quaternary moraine deposits.

These deposits are truncated in the northern and southern most extents by two E-W striking inward dipping normal faults, that dip between 70–80° to form a graben-like structure (Fig. 3b). The faults are identified as lithological discontinuities, mostly strikingly between the basaltic and andesitic lavas, and the red pyroclastic breccia (PadB). The faults can also be delimited, as the various units inside the valley can only be observed inside these boundaries and could not be correlated with any of the outcrops observed outside of the Quebrada Lunes valley. The lowermost extent of the graben faults can be further constrained since they do not cut the massive pyroclastic breccia (PB) at the base of the sequence. This unit instead, extends laterally, to the south, and downwards for about 4 km, seeming to overlay the basaltic and andesitic stratified lavas outside the boundaries of the valley.

The faults that are contained inside the valley are mostly normal faults, also forming graben-like structures and producing offset from 6 m up to 10 m, tilting some of the layers. There are at least nine faults cutting pyroclastic breccia (ApB and PadB) and tuff layers (PT and LaT2). Most of these displacements produced by the normal faults can be observed along the different gorges that make up the valley.

The deposits that outcrop inside the Quebrada Lunes Valley cannot be correlated with any of the surrounding rocks and in addition, almost 2 km north from the northernmost study area limit, specifically at 377.348 m E, 6.312.100 m N and 2.300 m.a.s.l., a plagioclase and quartz-bearing dacitic to rhyolitic tuff was found. This is the only pyroclastic unit found outside the valley at a similar topographic level and was similar in texture and composition to the LdT layer, thus it was possible to correlate.

[FIG.3]

4.2. Pyroclastic subvolcanic textures in Yerba Loca

In the entrance of the Yerba Loca national park (Fig. 2) we identify two subvolcanic rocks that cut the pyroclastic breccia that we describe at the bottom of the Quebrada Lunes Valley sequence (PB). In this area the pyroclastic breccia presents locally propylitic and phyllic alteration spatially related to subvolcanic rocks. These two units are (Fig. 4):

GpB. Pyroclastic breccia similar to PB but is lighter grey with an ash-rich matrix supported with polymict clasts up to 70 cm diameter. This breccia changes in colour to orange-light grey and a darker grey due to locally found hydrothermal alteration.

PD. Subvertical rhyo-dacitic pyroclastic dike, 5 m wide and striking NNE-SSW, which cut GpB formed by a competent central core surrounded by angular blocky monomict clasts, up to 20 cm diameter, supported in an incompetent ash matrix, similar to the core. Studies under the microscope indicate that this rock is formed by volcanic glass tuff with 0.5 to 1.5 mm lithic and glass fragments with a volcanic glass matrix, 0.6 to 1.5 mm crystals of plagioclase, quartz, biotite and amphibole, up to 3 mm juveniles and sharp Y and C-shaped shards in an ash matrix (Fig. 4d).

[FIG 4]

4.3. Composite volcano sequence at the Cerro Colorado base

At the north slope of the Molina River up to the Farellones village (Fig. 2), a 1.7 km thick sequence of units exposed throughout the Farellones valley. The lower part of the sequence is made up of more than 40 lava flows ranging in the thickness from 2 to 60 m, intercalated with block and ash flow (BAF; 10–35 m thick), a rhyolitic pyroclastic layer (RP; 15–25 m thick), and other pyroclastic deposits (OPD; 10–60 m thick), such as ash fall and tuffs, steeping 10 NE and cut by at least five leucocratic E-W dikes. Many, perhaps most, of the individual lava flows, exhibit evidence of autobrecciation at their bases. The explosive origin fall and flow deposits are recognised in aerial imagery by distinct changes in layer morphology and colour. For example, a distinctive orange unit, interpreted as a rhyolitic pyroclastic layer, presents a positive topographic expression and can be followed for several hundred metres throughout the valley (Fig. 5a, b).

The sequence continues in two blocks, separated by a NE-SW striking low angle thrust fault (Fig. 5c, d). In the western block (left side of Fig. 5c), we observed a 40 m thick andesitic lava flow; on top of this lava flow, we note a key unit, a 10 m thick greenish pyroclastic flow breccia (GB) overlain by rhyolitic pyroclastic deposits (RPD), with a minimum thickness of about 200 m. The upper contact of the RPD is not exposed so the thickness estimate is an underestimate. We are able to follow the green pyroclastic breccia (GB) for 1 km east, until it is cut by a thick sequence of explosive origin flow and fall deposits, corresponding to the eastern block (right side of Fig. 5c). In this block, at the base of the southernmost extent of the hanging wall of the reverse fault, we find a roughly 10 m sequence of well-exposed and easily accessible explosive origin fall and flow deposits (FFD). This deposit is overlain by rhyolitic pyroclastic deposits (RPD), with a minimum thickness of about 200 m, cut by a rhyolitic NW-SE dike (RD). The sequence ends with a thick layer of chaotically disposed rhyolites, which makes up the final > 300 m up to the Cerro Colorado hilltop, which is described in detail in the section called “Cerro Colorado hill Dome Complex”.

[FIG: 5]

We find a sequence of four pyroclastic flow and fall deposits (FFD, as mentioned previously) (Fig. 6). The lowest part of the sequence is a block and ash flow (BAF) with a polymict clast supported pyroclastic breccia and unwelded ash matrix. The clast sizes range between 4 cm up to very large blocks, greater than 1.25 m. The clasts and blocks have long axis/short axis ratios between 0.2 and 0.9, indicating that they have low sphericity and are mostly angular. The first part of the sequence has a thickness of 2.5 m, which is likely a minimum thickness since we were not able to trace the deposit laterally. A thin section from BAF matrix was made (Fig. 6d) and so, it is described as a lithic tuff with 44% of andesites and crystalline tuff fragments, 23% of sub-rounded and low sphericity scoria up to 4 mm juvenile clasts, and 19% of plagioclase sub-rounded fractured crystals and minor crystals of altered pyroxene in a (10%) vitric matrix.

Overlying and concordant with the block and ash flow was a 2.4 m thick layer of finely laminated volcanic ash. Lithic clasts are also present, but are much smaller, up to a maximum of about 10 cm. A thin section from these deposits was made (Fig. 6e, f), and from these thin sections, the deposits are described as a crystalline tuff with 27% of crystals between 1 and 3.8 mm of plagioclase, clinopyroxene and hornblende in a (73%) vitric matrix devitrificated to quartz. Overlying and concordant with the volcanic ash layer was a monomict breccia, a pyroclastic layer, which was deposited in a matrix of unlaminated volcanic ash, with angular and highly spherical clasts with diameters between 2 cm and 15 cm. And finally, the uppermost part of the sequence is a 1-meter minimum thickness layer of ash with monomict tuff clasts between 2 and 5 cm, and with coarse to fine lamination.

[FIG. 6]

4.4. Cerro Colorado hill Dome Complex

The Cerro Colorado hill represents a discrete topographic high with outcrops of a thick (> 300 m) package of sub-horizontal to steep rhyolitic chaotic lava flows (Fig. 7a). We note that the steepest rhyolitic lava flows contain spherulites. In general, the rhyolites contain 10–30% microfragmented crystals of plagioclase, hornblende, pyroxene, and quartz that range in size up to 3.8 mm, in 70–90% of vitric groundmass with minor crystals of quartz and plagioclase. They are made up from 6–50% of quartz and k-feldspar spherulites, with the lesser spherulites content in the sample from the top of the Cerro Colorado hill (Fig. 7b, c).

[FIG. 7]

In the area we find one rhyolitic dike with an attitude of N60W/60SW, emplaced within the rhyolitic pyroclastic deposits (RPD) as shown in Fig. 5 (RD). This dike is a rhyolitic porphyritic dike with 30% of phenocrystals of plagioclase and quartz in a vitric groundmass (70%); also, it has a content of 12% of quartz and k-feldspar spherulites (Fig. 8).

[FIG 8]

Due to the chaotic disposition of the rhyolitic lava flows, and the fractured crystals showed in the thin section, we interpret these rhyolites as a flow-banded dome complex, from the Cerro Colorado hill to the Farellones village, which transitions to a thick flow-banded rhyolitic lava flow.

Additionally, beyond the delimited study area, at least two rhyolitic domes with pyroclastic breccias separating the effusive facies is observed in El Pintor Hill, located 8 km north of the Cerro Colorado Dome (Fig. 9a). A flow-banded dome with abundant spherulites and holohyaline textures in the margin is exposed (Fig. 9b and 9c). This dome is intruded by rhyolitic and mafic dykes (beyond the scope of this study; Fig. 9a). The rhyolites are intercalated with green ash tuffs (GT) and pyroclastic breccias, predominantly monomict with felsic volcanic clast sub-angular with normal and inverse grading in a lapilli-rich matrix with minor ash fragments (Fig. 9d), interpreted as a block and ash flow (BAF). Nearby, at the top of the El Pintor rhyolites, fumarolic manifestations, are recorded as shown in Fig. 9e.

[FIG 9]

4.5. Estimation of the volume of volcanic products

An attempt to estimate the volume of volcanic deposits was made for the previously described products attributed to both lava flows, pyroclastic deposits and dome complexes within the specific study areas, with the intention to characterize the component of explosive and effusive volcanic activity in the zone. Three different calculations were made for the lava flows, pyroclastic deposits and dome complex from the Cerro Colorado sequence from the base to the top, and for the deposits related to the Quebrada Lunes and Yerba Loca pyroclastic products.

For the dome and pyroclastic deposits, related to the Cerro Colorado sequence; the minimum volume can be calculated from an interpolation of the dome, ash deposits, and pyroclastic deposits following the identified outcrops, and defining limits for the deposits from field observations. The north limit for these layers was defined at the Barros Negros gorge since the deposits were observed at the south slope of the gorge, but not on the north slope. The south limit was defined at the Molina river, since pyroclastic layers were not identified in the south slope. For the western limits, these are observed and interpreted as shown in Fig. 5. The eastern limits were inferred following the shape of the Cerro Colorado hill.

With these limits, an area was calculated using ArcGIS, and along with the thickness described in previous sections, a volume of the deposits could be calculated. For the pyroclastic deposits interbedded in the lava flows (RP, BAF1 and OPD; Fig. 5), the volume was calculated just by multiplying the area by the thickness, as if the layers were a block. For the pyroclastic deposits in the top part (RPD, GB and FFD; Fig. 5) and the Cerro Colorado Dome Complex, the volume was calculated as a third of the area multiplied by the thickness, as if these layers were a pyramid.

Since some minimum and maximum thickness were given to the layers, a range for the volume can be calculated, which represents the minimum volume due to the outcrops preserved nowadays from erosion. Furthermore, the pyroclastic deposits in the top (Ign, GB and FFD) are in both the footwall and the hanging wall of a thrust fault as interpreted in Fig. 5; and so, a substantial volume may have been eroded or displaced from the upper section of the unit.

Table 1

Estimated minimum volume for the explosive deposits related to the Cerro Colorado sequence.
Deposit	Cerro Colorado Dome Complex	Pyroclastic Deposits (RPD, GB, FFD)	Pyroclastic Deposits (RP, OPD, BAF1)
Thickness	300 m	590 m	35–120 m
Mapped Area	6.12 km²	8.07 km²	17.61 km²
Volume	0.613 km³	0.973 km³	0.617–2.114 km³

For the lava flows related to the Cerro Colorado sequence, the minimum volume can be estimated assuming all the subhorizontal lava flows found in the study area come from multiple effusive events erupted from the same Miocene volcanic centre. We estimate the volume assuming a block, with the same dimensions as the whole study area and a thickness of 0.7 km, which is between the beginning of the subhorizontal lavas up to the beginning of the rhyolitic lavas, part of the dome complex. This gives a total of 24.5 km³ from uneroded volcanic products, including both explosive and effusive deposits. From this, we subtract the already calculated minimum volume associated with the pyroclastic deposits, which comes to around 22 km³ of effusive deposits related to multiple lava flows. Taking into consideration the average thickness of 16 m for an individual lava flow, as this was possible to measure on the south side of the Cerro Colorado sequence, through drone pictures, we can also give a coarse estimate of the number of lava flows that made up the sequence, and the average volume of erupted material for a singular effusive eruption. This gives a lava flow volume of 0.5 km³ averaged over the 43 lava flows measured.

Lastly, for the pyroclastic deposits linked to the Quebrada Lunes Valley and Yerba Loca, the volume can be calculated as an inverted cone, as if the pyroclastic deposits were the result of explosive volcanism emplaced on top of the previous lava deposits, forming a cone shape structure, filled with pyroclastic deposits as the product of the eruption. This would be a 1.5 km diameter cone, with a height of at least 700 m, giving an approximate volume of 1.65 km³ of pyroclastic, from explosively derived material for the maar-diatreme volcanism.

5.1. Explosive maar-diatreme volcanism fed by pyroclastic dikes

Here we discuss the formation of a discrete maar-diatreme complex within the Quebrada Lunes valley. The following observations of units within the valley allow the interpretation of a maar-diatreme complex especially when they are combined with complementary observations from deeper in the diatreme complex around the Yerba Loca National Park. The main observations that support a maar-diatreme complex interpretation are as follows.

There is a tendency to form stratified deposits in the upper part and unstratified deposits in the lower parts of maar-diatreme complexes (Ross et al. 2017; McPhie et al. 1993). This is represented in the Quebrada Lunes sequence which exhibits stratified pyroclastic tuffs and breccias intercalated in the upper part of the valley, whereas the lower parts are characterized by a massive unstratified pyroclastic breccia (PB). The pyroclastic breccia (GpB) found at Yerba Loca and at the base of the Quebrada Lunes Valley sequence (PB), is interpreted as a subsurface diatreme breccia. These breccias are characterized by centimetre to meter sized matrix-supported heterogeneous clasts hosting hydrothermal fluids that in turn produce alteration (Ross et al. 2017). Maar-diatremes have a surface expression characterized by tuffs interbedded with pyroclastic breccias that may contain juvenile clasts (including shards) (Lorenz 1973; McPhie et al. 1993). The upper sequence of the Quebrada Lunes Valley hosts shard-bearing pyroclastic breccias interbedded with pyroclastic tuffs. Documented maar-diatreme systems produce a complex zone of brecciation in the form of steep crosscutting bodies filled with pyroclastic deposits (González & Greco 2019), similar to the pyroclastic dike described at the Yerba Loca entrance. This dike may represent a vestige of the conduit that fed the maar-diatreme volcanism (Díaz-Bravo & Morán-Centeno 2011).

Phreatomagmatic eruptions occur in clusters (Ross et al. 2017), giving place to phreatomagmatic breccias that cut prior phreatomagmatic breccias. This is relatable with the pyroclastic breccia (GpB) found at Yerba Loca, cutting a prior pyroclastic breccia (PB). The size of the Quebrada Lunes Valley as being around 1 km in diameter and around 700 m in depth fits well with described sizes for maar-diatreme craters in other parts of the world (e.g., Palladino et al. 2015; Valentine et al. 2011). Whilst an ejecta ring surrounding the maar-diatreme crater was not observed (White & Ross 2011), this is perhaps not surprising given the level of erosion in the area. Hints of a deposit that may comprise a part of the tuff ring could be represented in the zone by a plagioclase and quartz-bearing dacitic to rhyolitic tuff layer found almost 2 km north from the valley, but little other evidence exists.

Generally, the diatreme phreatomagmatic breccia is contained within the limits of the maar-diatreme structure, but the breccia does outcrop at the bottom of the Quebrada Lunes valley sequence and extends out of these limits. Moreover, this breccia does coincide with the lava and lava breccia deposits described outside the valley, nor with any deposit described in any other location. This implies that this breccia had to be deposited in a paleo-surface, escaping from the limits of the maar-diatreme structure which suggests that the collapse may have been on the flank of the volcano or a slope of the paleo topography permitting pyroclastic overflow to the west.

The maar-diatreme would then have followed a sequence of events to arrive at its final morphology, observed today as summarized in Fig. 10. In the pre-eruption stage, the area of the composite volcano was made up of a series of andesitic and basaltic lavas. The inner plumbing system of the volcano would have likely already hosted prior diatreme-type complexes and an active hydrothermal system. Meanwhile we did not find evidence of sedimentary units associated with the presence of a caldera lake or similar, nor do we note any features of the lava flows typical of magma-ice/water interaction, these options are not discarded, since the evidence may be found in places we could not access (e.g., Mee et al. 2006). Given there is no obvious evidence of standing water or ice, we favour a mechanism of explosive interaction with a shallow hydrothermal system as the likely driver of the phreatomagmatic explosion and maar formation (White 1996; Geshi & Oikawa 2008). Intrusive magma interaction with shallow hydrothermal waters is a known trigger of phreatomagmatic explosions (e.g., Geshi et al. 2016). This interpretation is further supported by the observation of a pyroclastic dyke lower in the sequence at Yerba Loca. Whilst we do not infer that this dike fed the Quebrada Lunes Maar, due to the lateral and vertical distance to this system. We suggest that similar dikes would have intruded an active hydrothermal system, generating explosive magma fragmentation to deep levels in the dike, and the formation of an explosion crater collapse at the surface. The pyroclastic dike found at Yerba Loca may represent the vestiges of the feeder conduit of the later maar-diatreme eruption.

In the syn-eruption stage, Le Corvec et al. (2018) studied similar diatreme system formation as being related to lateral magma diversion of a feeding system from an explosive eruption. Following this idea, the Quebrada Lunes maar-diatreme complex likely erupted on the flanks or outside of the central eruptive centre of a larger composite volcano. Magma feeding the composite volcano may then have experienced a diversion process triggered by an explosive eruption that resulted in the diatreme fragmentation, forming intra-diatreme sills as observed in the Quebrada Lunes sequence (LS). The described deposits (pyroclastic breccias and tuffs) that fill the maar-diatreme crater were likely produced in the syn-eruption stage along with the formation of an ejecta ring that surrounds the maar-diatreme crater, probably eroded posteriorly.

Maar-diatreme volcanism occurs in clusters and as parasitic systems in composite volcanoes (Ross et al. 2017), as such it is reasonable to assume the main eruption was followed by other similar or smaller eruptions in the zone, consistent with later, and smaller maar-diatreme structures in Yerba Loca, since the pyroclastic breccia found there (GpB), cuts the Quebrada Lunes pyroclastic breccia (PB). These following eruptions likely produced minor collapse on the Quebrada Lunes Maar, deforming and faulting the deposits inside. The diatreme system is not the only diatreme related environment in the zone. Piquer et al. (2015) mapped rhyolitic porphyry dikes and diatreme systems in the Rio Blanco-Los Bronces ore at 382.500 m E and 6.335.000 m N (WGS84 UTM 19S), 23 km north of the Quebrada Lunes maar-diatreme. This is the youngest intrusion in the Rio Blanco–San Francisco Batholith of late Miocene to early Pliocene age, culminating in the magmatism at ~ 4.69 Ma, which was fed by a deeper magma chamber (Piquer et al. 2015). So, we can assume that the Quebrada Lunes maar-diatreme formed at a similar time and with a similar process to those diatremes found in the Rio Blanco system. This indicates that there were multiple active volcanoes spaced within a few tens of kilometres.

Lastly, in the post-eruptive stage, intracrater faults could be formed due to later eruptions that provoke the deformation of recently formed pyroclastic deposits. There is no evidence for post-eruptive sediments filling the maar-diatreme crater at the top of the valley since the entire sequence was buried with post-eruptive volcanic products from other volcanic centres and moraines, forming the final morphology and outcrops observed.

[FIG 10]

5.2. The Cerro Colorado edifice forming sequence

The Cerro Colorado sequence, from the lava flows at the base up until the dome complex, is interpreted as a section of a composite volcano. The first phase of this volcano formed mainly lava flows, basaltic to andesitic in composition. These effusive eruptions were followed by interspersed-discrete explosive eruptions, generating thin layers of pyroclastic fall and flow deposits of predominantly rhyolitic composition. The second phase was predominantly explosive and was accompanied by sequences of edifice or dome collapse. This phase produced numerous fall and flow deposits, such as block and ash flows, and ash fall, pyroclastic density currents and pyroclastic breccias. This second phase is also characterized by the emplacement of the dome complex, ending the volcanic activity. The third and final phase is a coeval deformation phase represented by the dipping lava sequence and inverted faults within the Cerro Colorado hill.

The Cerro Colorado Dome is complex and can be broadly classified into three categories, with temporal and spatial transitions between the categories still poorly constrained. A section with an eroded carapace and absence of breccias on the upper surface dome which we classify as related to endogenous dome growth. An area with multiple smooth transitions from laterally constrained dome lavas to longer lava flows near the Farellones village which we characterize as Coulee related dome growth. An area with a high ratio of volcanic glass (obsidian) and spherulite content implying a devitrification process. The formation of domes can be linked to the presence of a series of rhyolitic dikes found at the top of the rhyolitic pyroclastic deposits (RPD; Fig. 5). As the domes and dikes possess similarities in composition and textural characteristics, represented by the partial formation of spherulites and fragmented crystals, it is reasonable to assume the dikes fed part of the dome complex.

As for the interpreted thrust fault (Fig. 5), we propose the possibility that this fault represents a reactivated and inverted normal fault, as is suggested for other faults in the same geological frame by other authors (e.g., Fock et al. 2006). Following this thread, considering a normal fault in the context of explosive style deposits and dome formation and collapse, we can also imply that this part of the sequence including pyroclastic deposits, normal (reactivated) faults and domes, can be attributed to either part of a larger caldera ring-fault system or reactivation of extensional accommodation structures. Since no further evidence to support the idea of a caldera is present in the Cerro Colorado hill, it remains speculative and as a suggestion for future study in the zone.

The late silicic volcanic episodes represented by the lithofacies observed in Fig. 9, may represent a widespread stage in and around the study area. For the dome exposed in the El Pintor Hill, we suggest exogenous construction of the dome fed partly by a rhyolitic dike that cuts the underlying deposits. Other mafic dikes are associated with later volcanic events (Fig. 9a). The field observations of the El Pintor Hill domes and interbedded pyroclastic breccias, suggest that the volcanic activity during this late stage was dominated by several discrete centres of rhyolitic composition, such as the El Colorado Dome, with intercalated explosive events, probably forming a volcanic field. These rhyolitic volcanoes probably hosted an extensive hydrothermal system, evidenced with fumarolic manifestations (Fig. 9e).

5.3. Size and location of the volcanic centres and eruptions

The calculated volumes for the explosive-related deposits can be attributed to at least three different medium size VEI 3 eruptions and one larger VEI 4 eruption using only the minimum estimated volume as the constraint. These VEI eruptions are comparable to prior eruptions of active Pliocene - Quaternary volcanoes, such as the Hudson caldera in the Aysén province of Chile, with a VEI 3–4 subplinian eruption on the 12th of August of 1991 (Naranjo et al. 1993) or the Puyehue-Cordon Caulle Volcanic complex in the Ranco and Osorno provinces of Chile, with a VEI 4 rhyolitic plinian eruption on the 4th of June of 2011.

The Cerro Colorado composite volcano (Fig. 11) proposed in this work relates to a previously proposed volcanic centre by Beccar (1986) and Rivano (1990), identified primarily as a topographic high. Since the area is highly eroded, the volcanic centre may in fact represent only a small portion, or a preliminary stage of volcanism that migrated north to form a larger volcanic complex. This is evidenced by numerous rhyolitic deposits and associated structures observed in the El Pintor hill to the north of the study area. Although, the rocks that form Cerro Colorado composite volcano correlates with the describe rocks from the upper member of the Farellones Formation indicated by Nyström et al. (2003), in this work we infer that the Quebrada Lunes maar-diatreme complex (Fig. 11) represents a posterior event, correlating in age to the upper member (or even later); due to its destructive nature. The main structure is located in Nystrom et al’s (2003) middle member and the pyroclastic breccia (PB) extends to the lower member. A final geologic map that summarizes the volcanic products and structures found in the area is shown in Fig. 11, naming the two proposed volcanic centres: Cerro Colorado composite volcano and the Quebrada Lunes Maar-Diatreme.

[FIG. 11]

Diversity in the active Pliocene-Quaternary volcanic arc, can also be observed in the Miocene volcanic centres, as expressions of numerous examples of both constructive and destructive volcanism, such as diatremes, dome growth and collapse events, ash and pyroclastic fall and flow deposits, extensive lava flows and internal intrusion and faulting episodes (Fig. 11). The findings in this work provide insight and a cross section of a flank derived maar-diatreme system, locally diverse deposits and structures, and destructive volcanism and processes related to the Andean Miocene activity. Our mapping distinguishes the main volcanic units and links to two main volcanic centres, complementing but building on prior maps that simply represented the rocks in the area as a single unit of three members, representative of the Farellones Formation.

Our analysis permits the identification of two Miocene volcanic centres in the study area.

The Quebrada Lunes Maar-Diatreme complex represents a collapse of about 1.5 km diameter and 700 m depth which produced high energy phreatomagmatic eruptions, characterized by phreatomagmatic, pyroclastic and collapse breccia deposits. It is proposed as a parasitic vent expressed as a flank collapse of the Cerro Colorado composite volcano erupting about 1.65 km³ of pyroclastic material.
The Cerro Colorado Composite Volcano represents a volcanic sequence centre that evolved from being constructed by a series of effusive dominant basaltic to andesitic lavas flows and with minor levels of pyroclastic forming explosive activity, to a dominantly explosive behaviour with the formation of at least 1.5 km³ of ignimbrites as well as ash-fall and block and ash flows terminating with a dominantly endogenous, obsidian derived coulée dome complex. The volcanism that formed this volcano, later migrated NE of this area, to the El Pintor Hill, to form a bigger volcanic system.

More detailed mapping needs to be completed in nearby areas to characterize other volcanic centres that gave rise to the Miocene volcanic lithologies from a physical volcanology perspective, indicating the type of volcano, style of eruption and related deposits. Such efforts will aid our understanding of the type and sequence of events potentially produced by active volcanoes in the area and also aid the potential for Geoheritage activities in the Miocene volcanics of the Chilean Andes.

Aguirre, L. (1960). Geología de los Andes de Chile central (provincia de Aconcagua): Santiago, Instituto de Investigaciones Geológicas Boletín, No. 9, 70 p. (with map at scale 1:100,000).
Aguirre, L., Feraud, G., Vergara, M., Carrasco, J., & Morata, D. (2000). 40Ar/39Ar ages of basic flows from the Valle Nevado stratified sequence (Farellones Formation), Andes of central Chile. In Proceedings, 9th Congreso Geológico Chileno, Puerto Varas (Vol. 1, pp. 583-585).
Amigo, A. (2021). Volcano monitoring and hazard assessments in Chile. Volcanica, 4(S1), pp.1-20, https://doi.org/10.30909/vol.04.S1.0120.
Beccar, I. (1983). Estudio geológico y petrográfico del área de Farellones—La Parva, Comuna de Las Condes, Región Metropolitana [Taller de título II]: Santiago, Universidad de Chile, Departamento de Geología, 49 p.
Beccar, I., Vergara, M., & Munizaga, F. (1986). Edades K-Ar de la Formación Farellones en el cordón del Cerro la Parva, Cordillera de los Andes de Santiago, Chile.
Bertín, D., & Orozco, G. (2015). Petrografía, geoquímica y peligros volcánicos de los volcanes Palomo y Tinguiririca, VI Región. In Congreso Geológico Chileno (No. 14, pp. 124-127).
Bertin, D., & Silva, C. (2015). Geología y peligros del volcán Tupungatito, Región Metropolitana de Santiago. In XVI Congreso Geológico Chileno.
Charrier, R., Baeza, O., Elgueta, S., Flynn, J. J., Gans, P., Kay, S. M., ... & Zurita, E. (2002). Evidence for Cenozoic extensional basin development and tectonic inversion south of the flat-slab segment, southern Central Andes, Chile (33–36 SL). Journal of South American Earth Sciences, 15(1), 117-139, https://doi.org/10.1016/S0895-9811(02)00009-3.
Charrier, R., Ramos, V. A., Tapia, F., & Sagripanti, L. (2015). Tectono-stratigraphic evolution of the Andean Orogen between 31 and 37 S (Chile and Western Argentina). Geological Society, London, Special Publications, 399(1), 13-61.
Deckart, K., Clark, A. H., Aguilar, C., Vargas, R., Bertens, A., Mortensen, J. K., and Fanning, M. (2005). Magmatic and hydrothermal chronology of the giant Rio Blanco porphyry copper deposit, central Chile: Implications of an integrated U-Pb and Ar- 40/Ar-39 database: Economic Geology, v. 100, p. 905-934, https://doi.org/10.2113/gsecongeo.100.5.905.
Deckart, K., Godoy, E., Bertens, A., Jerez, D., & Saeed, A. (2010). Barren Miocene granitoids in the Central Andean metallogenic belt, Chile: Geochemistry and Nd-Hf and U-Pb isotope systematics. Andean Geology, 37(1), 1-31.
Diaz, E. F. G. (1961). La petrografía del cerro Tupungato y de otras rocas efusivas de la región (provincia de Mendoza). Revista de la Asociación Geológica Argentina, 16(3-4), 205-234.
Díaz-Bravo, B. A., & Morán-Zenteno, D. J. (2011). The exhumed Eocene Sultepec-Goleta Volcanic Center of southern Mexico: Record of partial collapse and ignimbritic volcanism fed by wide pyroclastic dike complexes. Bulletin of Volcanology, 73, 917-932, https://doi.org/10.1007/s00445-011-0460-5.
Drake, R. E., Curtis, G., & Vergara, M. (1976). Potassium-argon dating of igneous activity in the central Chilean Andes—latitude 33° S. Journal of Volcanology and Geothermal Research, 1(3), 285-295, https://doi.org/10.1016/0377-0273(76)90012-3.
Elgueta, S., Charrier, R., Aguirre, R., Kieffer, G., & Vatin-Perignon, N. (1999). Volcanogenic sedimentation model for the Miocene Farellones Formation, Andean Cordillera, central Chile. In Proceedings of the IV International Symposium on Andean Geodynamics, Göttingen.
Farías, M., Charrier, R., Carretier, S., Martinod, J., Fock, A., Campbell, D., ... & Comte, D. (2008). Late Miocene high and rapid surface uplift and its erosional response in the Andes of central Chile (33–35 S). Tectonics, 27(1).
Fock, A., Charrier, R., Farías, M., & Muñoz, M. (2006). Fallas de vergencia oeste en la Cordillera Principal de Chile Central: Inversión de la cuenca de Abanico (33-34 S). Revista de la Asociación Geológica Argentina, Publicación Especial, 6, 48-55.
Geshi, N. and Oikawa, T. (2008). Phreatomagmatic eruptions associated with the caldera collapse during the Miyakejima 2000 eruption, Japan. Journal of volcanology and geothermal research, 176(4), pp.457-468, https://doi.org/10.1016/j.jvolgeores.2008.04.013.
Geshi, N., Iguchi, M. and Shinohara, H. (2016). Phreatomagmatic eruptions of 2014 and 2015 in Kuchinoerabujima Volcano triggered by a shallow intrusion of magma. Journal of Natural Disaster Science, 37(2), pp.67-78, https://doi.org/10.2328/jnds.37.67.
Giambiagi, L. B., Ramos, V. A., Godoy, E., Alvarez, P. P., & Orts, S. (2003). Cenozoic deformation and tectonic style of the Andes, between 33 and 34 south latitude. Tectonics, 22(4).
Godoy, E. (2012). Sobre el variable marco geotectónico de las Formaciones Abanico y Farallones y sus equivalentes al sur de los 35° LS. Revista de la Asociación Geológica Argentina, 69(4), 570-577.
Godoy, E., Yañez, G., & Vera, E. (1999). Inversion of an Oligocene volcano-tectonic basin and uplifting of its superimposed Miocene magmatic arc in the Chilean Central Andes: first seismic and gravity evidences. Tectonophysics, 306(2), 217–236. doi:10.1016/s0040-1951(99)00046-3, https://doi.org/10.1016/S0040-1951(99)00046-3.
González, S. N., & Greco, G. A. (2019). Diques piroclásticos en el Complejo Volcánico Marifil. In XIII CONGRESO DE MINERALOGÍA, PETROLOGÍA ÍGNEA Y METAMÓRFICA, Y METALOGÉNESIS.
Gonzalez-Ferrán, O. (1994). Volcanes de Chile. Instituto Geográfico Militar, 640 p.
Instituto de Investigaciones Geológicas (Santiago, Chile), Thiele, R., & Cubillos, E. (1980). Hoja Santiago: región metropolitana: carta geológica de Chile escala 1: 250.000. Instituto de Investigaciones Geológicas.
Jara, P., & Charrier, R. (2014). Nuevos antecedentes estratigráficos y geocronológicos para el Meso-Cenozoifgco de la Cordillera Principal de Chile entre 32 y 32 30'S: Implicancias estructurales y paleogeográficas. Andean geology, 41(1), 174-209, http://dx.doi.org/10.5027/andgeoV41n1-a07.
Jara, P., Likerman, J., Charrier, R., Herrera, S., Pinto, L., Villarroel, M., & Winocur, D. (2018). Closure type effects on the structural pattern of an inverted extensional basin of variable width: Results from analogue models. Journal of South American Earth Sciences, 87, 157-173.
Jara, P., Likerman, J., Winocur, D., Ghiglione, M. C., Cristallini, E. O., Pinto, L., & Charrier, R. (2015). Role of basin width variation in tectonic inversion: insight from analogue modelling and implications for the tectonic inversion of the Abanico Basin, 32–34 S, Central Andes. Geological Society, London, Special Publications, 399(1), 83-107.
Klohn, C. (1960). Geología de la Cordillera de los Andes de Chile central, Instituto de Investigaciones Geológicas. Boletín, 8.
Le Corvec, N., Muirhead, J. D., & White, J. D. (2018). Shallow magma diversions during explosive diatreme-forming eruptions. Nature communications, 9(1), 1459, https://doi.org/10.1038/s41467-018-03865-x.
Lefebvre, N. S., White, J. D. L., & Kjarsgaard, B. A. (2013). Unbedded diatreme deposits reveal maar-diatreme-forming eruptive processes: Standing Rocks West, Hopi Buttes, Navajo Nation, USA. Bulletin of Volcanology, 75, 1-17.
Lorenz, V. (1973). On the formation of maars. Bulletin volcanologique, 37(2), 183-204, https://doi.org/10.1007/BF02597130.
McPhie, J. (1993). Volcanic textures: a guide to the interpretation of textures in volcanic rocks.
Mee, K., Tuffen, H. and Gilbert, J.S. (2006). Snow-contact volcanic facies and their use in determining past eruptive environments at Nevados de Chillán volcano, Chile. Bulletin of volcanology, 68(4), pp.363-376, https://doi.org/10.1007/s00445-005-0017-6.
Michol, K. A., Russell, J. K., & Andrews, G. D. M. (2008). Welded block and ash flow deposits from Mount Meager, British Columbia, Canada. Journal of Volcanology and geothermal research, 169(3-4), 121-144, https://doi.org/10.1016/j.jvolgeores.2007.08.010.
Mosolf, J. G. (2013). Stratigraphy, structure, and geochronology of the Abanico Formation in the Principal Cordillera, central Chile: Evidence of protracted volcanism and implications for Andean tectonics. University of California, Santa Barbara.
Muñoz-Gómez, M., Fuentes, F., Vergara, M., Aguirr, L., Olov Nyström, J., Féraud, G., & Demant, A. (2006). Abanico East Formation: petrology and geochemistry of volcanic rocks behind the Cenozoic arc front in the Andean Cordillera, central Chile (33° 50'S). Revista geológica de Chile, 33(1), 109-140, http://dx.doi.org/10.5027/andgeoV33n1-a05.
Naranjo, J. A., Moreno, H., & Banks, N. G. (1993). La erupción del volcán Hudson en 1991 (46° S): Región IX, Aisén, Chile.
Nyström, J. O., Vergara, M., Morata, D., & Levi, B. (2003). Tertiary volcanism during extension in the Andean foothills of central Chile (33 15′–33 45′ S). Geological Society of America Bulletin, 115(12), 1523-1537, https://doi.org/10.1130/B25099.1.
Palladino, D. M., Valentine, G. A., Sottili, G., & Taddeucci, J. (2015). Maars to calderas: end-members on a spectrum of explosive volcanic depressions. Frontiers in Earth Science, 3, 36, https://doi.org/10.3389/feart.2015.00036.
Pilger, R. H. (1984). Cenozoic plate kinematics, subduction and magmatism: South American Andes. Journal of the Geological Society, 141(5), 793-802, https://doi.org/10.1144/gsjgs.141.5.0793.
Pineda, C., Hammer, J., First, E. and Morata, D. (2021). Storage conditions of a caldera-forming volcanic eruption: Insights from the Pudahuel rhyolitic ignimbrite in central Chile (32° 10′ S). Lithos, 400, p.106382.
Pinto, L., Muñoz, C., Nalpas, T., & Charrier, R. (2010). Role of sedimentation during basin inversion in analogue modelling. Journal of Structural Geology, 32(4), 554-565.
Piquer, J. (2015). Structural geology of the Andes of central Chile: controls on magmatism and the emplacement of giant ore deposits (Doctoral dissertation, University of Tasmania).
Piquer, J., Hollings, P., Rivera, O., Cooke, D. R., Baker, M., & Testa, F. (2017). Along-strike segmentation of the Abanico Basin, central Chile: new chronological, geochemical and structural constraints. Lithos, 268, 174-197, https://doi.org/10.1016/j.lithos.2016.10.025.
Piquer, J., Skarmeta, J., & Cooke, D. R. (2015). Structural evolution of the Rio Blanco-Los Bronces District, Andes of Central Chile: controls on stratigraphy, magmatism, and mineralization. Economic Geology, 110(8), 1995-2023, http://dx.doi.org/10.2113/econgeo.110.8.1995.
Rivano, S., Godoy, E., Vergara, M., & Villarroel, R. (1990). Redefinición de la Formación Farellones en la Cordillera de los Andes de Chile Central (32-34 S). Andean Geology, 17(2), 205-214.
Ross, P. S., Carrasco Núñez, G., & Hayman, P. (2017). Felsic maar-diatreme volcanoes: a review. Bulletin of Volcanology, 79(2), 1-33, http://dx.doi.org/10.1007/s00445-016-1097-1.
SERNAGEOMIN. (2003). Mapa Geológico de Chile: versión digital. Servicio Nacional de Geología y Minería, Publicación Geológica Digital, No. 4 (CD-ROM, versión1.0, 2003). Santiago.
Serrano, L., Vargas, R., Stambuk, V., Aguilar, C., Galeb, M., Holmgren, C., ... & Stern, C. R. (1998). The late Miocene to early Pliocene Río Blanco-Los Bronces copper deposit, central Chilean Andes https://doi.org/10.5382/SP.05.09,.
Sruoga, P., Llambías, E. J., Fauqué, L., Schonwandt, D., & Repol, D. G. (2005). Volcanological and geochemical evolution of the Diamante Caldera–Maipo volcano complex in the southern Andes of Argentina (34° 10′ S). Journal of South American Earth Sciences, 19(4), 399-414, https://doi.org/10.1016/j.jsames.2005.06.003.
Stern, C. R., & Skewes, M. A. (1995). Miocene to present magmatic evolution at the northern end of the Andean Southern Volcanic Zone, Central Chile. Andean Geology, 22(2), 261-272.
Thiele, R., Beccar, I., Levi, B., Nyström, J. O., & Vergara, M. (1991). Tertiary Andean volcanism in a caldera-graben setting. Geologische Rundschau, 80(1), 179-186, https://doi.org/10.1007/BF01828775.
Thomas, H. (1953). Informe de la comisión geológica Thomas-Junge sobre la alta cordillera entre el rio Aconcagua y el rio Colorado. CORFO, Santiago, Chile.
Valentine, G. A., Shufelt, N. L., & Hintz, A. R. (2011). Models of maar volcanoes, Lunar crater (Nevada, USA). Bulletin of Volcanology, 73, 753-765.
Vergara, M., Charrier, R., Munizaga, F., Rivano, S., Sepulveda, P., Thiele, R., & Drake, R. (1988). Miocene volcanism in the central Chilean Andes (31 30′ S–34 35′ S). Journal of South American Earth Sciences, 1(2), 199-209, https://doi.org/10.1016/0895-9811(88)90038-7.
White, J. D., & Ross, P. S. (2011). Maar-diatreme volcanoes: A review. Journal of Volcanology and Geothermal Research, 201(1-4), 1-29, https://doi.org/10.1016/j.jvolgeores.2011.01.010.
White, J.D., (1996). Impure coolants and interaction dynamics of phreatomagmatic eruptions. Journal of Volcanology and Geothermal Research, 74(3-4), pp.155-170, https://doi.org/10.1016/S0377-0273(96)00061-3.

Supportinginformation.pdf

Download PDF

Version 1

posted

You are reading this latest preprint version

Diversity of volcanism and evidence of discrete eruption centres in the Miocene Andes of Central Chile

Status:

Version 1

Abstract

Figures

1. Introduction

2. Geological setting

3. Methods

4. Results

4.1. Evidence of maar-diatreme volcanism in Quebrada Lunes Valley

4.2. Pyroclastic subvolcanic textures in Yerba Loca

4.3. Composite volcano sequence at the Cerro Colorado base

4.4. Cerro Colorado hill Dome Complex

4.5. Estimation of the volume of volcanic products

5. Discussion

5.1. Explosive maar-diatreme volcanism fed by pyroclastic dikes

5.2. The Cerro Colorado edifice forming sequence

5.3. Size and location of the volcanic centres and eruptions

6. Conclusions

References

Supplementary Files

Status:

Version 1